The Neogene and Quaternary of England: landscape evolution, tectonics, climate change and their expression in the geological record

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The Neogene and Quaternary of England: landscape evolution, tectonics, climate change and their expression in the geological record Jonathan R. Leea,* , Ian Candyb , Richard Haslama a b

British Geological Survey, Keyworth, Nottingham, NG12 5GG, UK Department of Geography, Royal Holloway University of London, Egham, Surrey, TW20 0EX, UK

A R T I C L E I N F O

Article history: Received 18 May 2017 Received in revised form 16 October 2017 Accepted 24 October 2017 Available online xxx Keywords: Neogene Quaternary Landscape evolution Climate forcing Tectonics Alpine Orogeny Denudational isostasy

A B S T R A C T

During the Neogene and Quaternary, tectonic and climatic processes have had a profound impact upon landscape evolution in England and, perhaps as far back as 0.9 Ma, patterns of early human occupation. Until the Late Miocene, large-scale plate tectonic processes were the principal drivers of landscape evolution causing localised basin inversion and widespread exhumation. This drove, in places, the erosion of several kilometres of Mesozoic cover rocks and the development of a regional unconformity across England and the North Sea Basin. By the Pliocene, the relative influence of tectonics on landscape evolution waned as the background tectonic stress regime evolved and climatic influences became more prominent. Global-scale climate-forcing increased step-wise during the Plio-Pleistocene amplifying erosional and depositional processes that operated within the landscape. These processes caused differential unloading (uplift) and loading (subsidence) of the crust (‘denudational isostasy’) in areas undergoing net erosion (upland areas and slopes) and deposition (basins). Denudational isostasy amplified during the Mid-Pleistocene Transition (c.0.9 Ma) as landscapes become progressively synchronised to large-scale 100 ka ‘eccentricity’ climate forcing. Over the past 0.5 Ma, this has led to the establishment of a robust climate record of individual glacial/interglacial cycles enabling comparison to other regional and global records. During the Last Glacial-Interglacial Transition and early Holocene (c.16–7 ka), evidence for more abrupt (millennial/centennial) scale climatic events has been discovered. This indicates that superimposed upon the longer-term pattern of landscape evolution is a more dynamic response of the landscape to local and regional drivers. © 2017 Natural Environment Research Council, as represented by the British Geological Survey. Published by Elsevier Ltd on behalf of The Geologists' Association. All rights reserved.

1. Introduction Traditionally, the Late Cenozoic (Neogene and Quaternary) has been considered a somewhat benign period of UK Earth History with comparatively little landscape development occurring over several tens of millions of years (e.g. Hancock and Rawson, 2002). Relative to past global-scale tectonic events (e.g. Caledonian and Variscan Orogenic phases) this assessment perhaps holds some truth. Nevertheless, continuing from the Late Mesozoic and Palaeogene, the landscape from the beginning of the Neogene Period (23.02 Ma) has been in a state of constant flux, with the form of the modern landscape owing its existence to a wide range of tectonic and climatic drivers that have controlled geological processes (Westaway et al., 2009; Candy et al., 2010; Green et al.,

* Corresponding author. E-mail address: [email protected] (J.R. Lee).

2012; Gibbard and Lewin, 2003, 2016; Westaway, 2017). By 23 Ma, global climate had cooled progressively from the ‘greenhouse climates’ that dominated earlier parts of the Palaeogene (Zachos et al., 2001, 2008) to the onset of ‘icehouse climates’ and the onset of large-scale glaciation in Antarctica (Kennett, 1977; Zachos et al., 1992) (Fig. 1a). The global configuration of the continents also broadly resembled the contemporary palaeogeography, with continued opening of the North Atlantic and closure of the Tethys Ocean leading to the formation of the Alpide Mountain Belt across southern and central Europe and Asia (Fig. 1b). Ongoing tectonics, corresponding to Alpine compression and the effects of the Iceland Mantle Plume strongly influenced landscape evolution during the preceding Palaeogene (Gibbard and Lewin, 2003; Newell, 2014) continuing into the Early and Mid-Miocene (Tiley et al., 2004; Williams et al., 2005; Westaway, 2010). However, during the Late Miocene the impact of these regimes subsided as the regional tectonic stress regime evolved. Instead, landscape evolution was driven by differences in relative crustal loading caused by vertical

https://doi.org/10.1016/j.pgeola.2017.10.003 0016-7878/© 2017 Natural Environment Research Council, as represented by the British Geological Survey. Published by Elsevier Ltd on behalf of The Geologists' Association. All rights reserved.

Please cite this article in press as: J.R. Lee, et al., The Neogene and Quaternary of England: landscape evolution, tectonics, climate change and their expression in the geological record, Proc. Geol. Assoc. (2017), https://doi.org/10.1016/j.pgeola.2017.10.003

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Fig. 1. (a) The Cenozoic time-scale showing major climatic trends (from Zachos et al., 2008; Newell, 2014) with key tectonic (brown) and climatic (warm – red, cold – blue) events. Global palaeogeography during the Mid-Miocene (b) and following the last Quaternary glaciation (c). Based on maps produced by C.R. Scotese (2002); http://www. scotese.com, (PALEOMAP website).

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uplift, erosion and sedimentation (‘denudational isostasy’) (Watts et al., 2000, 2005; Westaway et al., 2002; Lane et al., 2008; Westaway, 2017). This driver of landscape change has persisted through to the present day albeit against a backdrop of progressive climate cooling and increased seasonality, which has amplified denudation rates. It culminated at the beginning of the Quaternary (2.588 Ma) with the establishment of glaciation in upland areas bordering the North Atlantic (Flesche Kleiven et al., 2002; Sejrup et al., 2005; Knies et al., 2009; Lee et al., 2012; Reinardy et al., 2017) and development of regular short-term cyclical changes in sealevel, glaciation and climate (Rose, 2010) (Fig. 1c). Understanding both the longer-term and shorter-term influences of tectonics and climate during the Late Cenozoic is crucial to understanding the nature and timescales of landscape change and the nature and properties of the shallow sub-surface (the ‘zone of human-interaction’). In this paper, we review the evidence for landscape evolution during and since the Neogene, particularly within the broader context of regional/global tectonic and climatic processes. One notable omission from the paper is the role of modern humans as major drivers of landscape change and this topic is discussed elsewhere in this Special Issue (Zalasiewicz et al., this volume). 2. Drivers of Neogene and Quaternary landscape change This section of the paper aims to provide a broad context to the two principal mechanisms – tectonics and climate, which have driven landscape change in England during the Neogene and Quaternary. 2.1. Tectonics Until recently, England, the wider UK and northeast Atlantic margin has been considered largely passive (tectonically) during the Cenozoic. However, despite residing within an intracratonic setting along the northwest part of the Eurasian plate, England and adjacent onshore and offshore regions have been tectonically active during and following the Palaeogene (Hillis et al., 2008; Westaway, 2017). Indeed, during the Neogene and Quaternary tectonics have played a significant long-term role in controlling the surface geology of England, the shape and form of the landscape as well as the development of major hydrocarbon resources both onshore and offshore (Hillis et al., 2008; Holford et al., 2009; Westaway, 2017). In the broader regional tectonic context, many of the major Mesozoic basins of north-western Europe started to undergo exhumation (uplift and erosion) during the Cretaceous and this continued through the Cenozoic with progressive subsidence occurring in adjacent marginal areas. Subsidence has led to the development of the European Cenozoic Rift System which extends northwards beneath part of the North Sea Basin (Ziegler, 1992; Kockel, 2002; Schumacher, 2002; Dèzes et al., 2004) forming a major Neogene and Quaternary depositional centre (Jordt et al., 1995; Kuhlmann et al., 2006a). Other active Cenozoic basins that are relevant to the landscape of England include the Irish Sea Basin and English Channel although these have only become major depositional centres during the Late Quaternary. In some areas of England, greater than 2 km of Cenozoic exhumation has occurred with a significant component occurring since the beginning of the Miocene (Hillis et al., 2008). Eight different major exhumation events have affected the UK between 120 Ma (Early Cretaceous) and 15 Ma (Miocene). Two exhumations events occurred during the Miocene (Holford et al., 2009) with three additional phases recognised during the Pliocene and Pleistocene (Watts et al., 2000, 2005; Westaway et al., 2002; Lane et al., 2008; Westaway, 2009b). Exhumation formed several major tectonic unconformities that

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can be mapped onshore across parts of the UK and extending offshore through the North Sea and the Northwest European Atlantic Margin and into continental Europe (Stoker et al., 2005; Holford et al., 2009). Three long-term and one short-term tectonic mechanisms have operated during the Late Cenozoic and influenced landscape development. North Atlantic Rifting commenced during the Mesozoic and resulted in the break-up of the Laurasia supercontinent adjacent to what is now the UK during the Early Cretaceous to Palaeogene (Ziegler, 1988; Doré et al., 1999). This ‘ridge-push’ caused progressive basin-wards tilting of the continental margins including the eastwards tilting of central and eastern England towards the North Sea. Also linked to the opening of the North Atlantic was, during the Palaeogene, an intense period of magmatism associated with the Iceland Mantle Plume which generated the British Tertiary Igneous Provenance in northern Britain (Mussett et al., 1988; Hansen et al., 2009). Plume activity also caused crustal thickening (magmatic underplating) (Brodie and White, 1994; Thomson, 1995; Jones et al., 2002) and widespread exhumation across western and northern Britain (Tiley et al., 2004; Williams et al., 2005; Westaway, 2010) resulting in the removal of younger Mesozoic ‘cover’ (Green et al., 2012). Plume-related uplift is considered to have occurred over twodifferent temporal scales. Firstly, as a prolonged phase of uplift driven by thermal swell and magamtic underplating that occurred through much of the Palaeocene-Eocene (Brodie and White, 1994; Jones et al., 2012). Secondly, shorter transient pulses of uplift on the margins of the thermal swell each lasting about a million years associated with ‘hot-blob convection’ within the upper asthenosphere (Shaw Champion et al., 2008; Jones et al., 2012). The spatial influence of plume-related uplift is a matter of some debate. Previously, uplift has been linked to areas more proximal to plume activity such as western and northern Britain. However, within another paper in this Special Issue, Gale and Lovell (this volume) have argued that prolonged plume-related uplift drove erosion of upto 500 m of chalk from the Chilterns-East Anglia dome and caused the southeastwards tilting of England. Alpine Compression was initiated in the UK during the Late Cretaceous due to the closure of the Tethys Ocean that once separated the ancient continental landmasses of Gondwana and Laurasia. In Europe, collision of the Eurasian-Iberian and African tectonic plates led to a major phase of mountain building with the generation of the Alpide belt that include several major western mountain ranges (e.g. the Alps and Pyrenees) that extend across central and southern Europe (Ziegler et al., 1995; Cloetingh et al., 2005; Hillis et al., 2008). Although this collision zone was situated over 1000 km to the south of England, northwards directed compressive stresses have nevertheless been transmitted into the interior of the Eurasian Plate (Ziegler et al., 1995), deforming much of the UK and the northeast Atlantic margin (Stoker et al., 2005). Within the Wessex-Weald Basin, for example, inversion during the earlier Pyrenean orogenic phase (Eocene-Oligocene) led to localised exhumation over major geological structures such as the Weald-Artois Anticline and the Portland-Wight disturbance (Ziegler, 1990; Ziegler and Dèzes, 2007). A later (Eocene-Miocene) tectonic phase (the Savian) caused further regional uplift and basin inversion affecting the Sole Pit Trough, Flamborough Head fault zone and Cleveland Basin (Kirby and Swallow, 1987; Van Hoorn, 1987). Denudational (erosional) isostasy occurs where the transfer of an applied crustal load from upland to basinal areas drives upland uplift (unloading) and basinal subsidence (loading) (Gilchrist and Summerfield, 1991; Bishop, 2007). To quantify this, in areas of isostatic equilibrium, studies have shown that approximately 0.85 km of surface uplift occurs in response to the removal of a kilometre of crustal load (Bishop, 2007). During the Late Neogene

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and Quaternary, the significance of this mechanism has amplified. This is because the enhancement of climatic seasonality has accelerated both weathering and erosion rates, and improved the efficiency of surface processes at recycling sediment from upland source areas to basinal sinks (see Section 2.2). Despite this, there is considerable debate within the literature whether climate change is the cause or effect of denudational isostasy (see Molnar and England, 1990; Bishop, 2007 for discussion). A final shorter-term tectonic influence on landscape development is that of glacial isostatic adjustment. Stresses imposed by large bodies of ice progressively overprint the prevailing tectonic stress regime causing deformation within both the crust and upper mantle. This can occur through vertical glacier loading and the development of a forebulge in front of the glacier as the crust and mantle are laterally displaced (Stewart et al., 2000). The ground surface beneath the thickest ice load can be depressed by upto 500 m with uplift in the forebulge area being lower and in the range of a few metres to several hundred metres (Lambeck, 1995; Stewart et al., 2000). During and following deglaciation, elastic rebound of the crust occurs as the ice load is removed. In addition, retraction of the forebulge and the development of a dome of crustal extension (which expands outwards from the centre of the former ice mass) occurs when the displaced mantle flows back and the horizontal stresses relax. The duration of isostatic adjustment is difficult to reconcile although studies in Fennoscandia suggest that current crustal recovery following the last deglaciation may last for 40,000 years (Whitehouse, 2006). In Britain, rates of isostatic uplift are generally largest in northern England and Scotland where the greatest volume of crustal ice loading has been removed; in contrast, much of southern and central England which lay at or beyond the limits of Late Devensian glaciation are either in isostatic equilibrium or undergoing gentle subsidence (Shennan and Horton, 2002; Shennan et al., 2006; Bradley et al., 2011). Glacial isostatic adjustment is also likely to result in an increase in seismicity, fault reactivation and crustal fracturing (Firth and Stewart, 2000; Stewart et al., 2000; Lagerbäck and Sundh, 2008; Pascal et al., 2010). This is especially true during deglaciation when strain stored within the crust is released as the vertical load is removed. During ice loading, by contrast, the application of a vertical stress to compressional and strike-slip tectonic regimes cause the differential stress in the crust to be reduced, resulting in faults locking and stabilising (Johnston, 1987). The relative influence of each of the four tectonic mechanisms on landscape evolution has evolved through the Cenozoic and this is the principal focus of the tectonic element of this paper. For the past ten million years, England has been subjected to a low-stress compressional to intermediate tectonic stress regime. This implies that only weakly compressional and strike-slip systems will have been tectonically active. Extensional tectonic systems, by contrast, will only have become activated in response to changes in the applied vertical load (see Discussion). It is also relevant to point out that there is significant debate on how isostasy is accommodated within the lithosphere. Current models suggest that lithospheric isostatic response occurs through a combination of flexural isostasy (elastic flexure) and flexural rigidity (bending of nonrigid materials) (see Allen and Allen, 2013 for overview). Key factors driving lithospheric response include the rigidity of the lithosphere undergoing flexure (a function of composition, temperature and density) and the shape and wavelength of the applied load. For instance, larger wavelength lithospheric loads such as major ocean basins are generally fully-compensated by flexural isostasy. By contrast smaller wavelength lithospheric loads, including individual mountains and valleys, achieve isostatic compensation by flexural rigidity. A key and highly-controversial

issue relates to the precise scale at which flexural isostasy and flexural rigidity occur and critically how they can be recognised within the geological record. 2.2. Climate Following the Palaeocene-Eocene Thermal Maximum (PETM; c. 55.5 Ma) global climate began to cool progressively through the remainder of the Palaeogene and following Neogene and Quaternary periods (Zachos et al., 2001, 2008; Lisiecki and Raymo, 2007; Fig. 1a). The inception of large-scale glaciation occurred first in Antarctica (c. 34 Ma) and then in the Northern Hemisphere, Greenland (c. 12.7 Ma), during the Miocene (Fig. 1a). Within the Pliocene and Pleistocene, orbital forcing was the key driver of climate change which in-turn has influenced Earth surface processes and landscape response. Orbital forcing is a term used to describe changes in the shape of the Earth’s orbit around the sun that drive cyclic variations in the seasonal distribution of insolation (solar radiation) across the Earth’s surface. These orbital cycles, called Milankovitch Cycles, have occurred throughout geological time and are also recognised during some older parts of the British geological record (e.g. Leeder, 1988; Mutterlose and Ruffell, 1999; Bonis et al., 2010). However, during the Plio-Pleistocene they acted in tandem with global-scale geography and oceanic circulation patterns to accentuate the global climate signal. During the Neogene and Quaternary, orbital variations referred to as eccentricity (100 ka), obliquity (40 ka) and precession (20 ka) cycles, are widely acknowledged to have been the principal driver of Ice Ages and generate glacial (cold)/interglacial (temperate) climate cycles. These cycles are expressed in the geological record of England and are important agents of landscape change. However, it is important to highlight that by itself, orbital forcing does not generate sufficient change in solar insolation to explain the magnitude of the temperature variations encountered. Instead, changes in solar insolation have been amplified by factors such as albedo effect and greenhouse gas uptake/release to enhance the climate forcing. Long-term Cenozoic climate trends are reconstructed from foraminfera assemblages and sea-surface temperatures obtained from marine sediment sequences (Zachos et al., 2008). The resulting marine isotope record, based upon the benthic d18O record from multiple marine sequences, is then statistically stacked and smoothed to produce a single record. For the Quaternary a near-complete record has been produced and is routinely employed as the global climate signal, with sub-division into individual Marine Isotope Stages (MIS) back through geological time – odd numbers representing warm (interglacial) stages and even numbers corresponding to cold (glacial) stages (Lisiecki and Raymo, 2007). The current paradigm within British Quaternary studies is to use a combination of relative stratigraphic and absolute-dating techniques to produce a regional stratigraphy that can be correlated with distinct marine isotope stages. The magnitude and frequency of climate cyclicity has evolved over the past 4 Ma with a trend towards progressively longer and more extreme periods of global cooling. The Pliocene (Late Neogene) is primarily characterised by 20 ka (precession-driven) cycles that, as well as having a high-frequency, are characterised by limited cooling/ice mass development. The onset of the Quaternary (2.588 Ma) saw the increased dominance of the 40 ka (obliquitydriven) cycles. These were characterised by a shift towards greater degrees of cooling but, using the benthic d18O signal as a benchmark, ice volume build-up during glacials that would have been comparable, or even less than the ice volume increases that occurred during the cooler sub-stages of warm marine isotope stages during the Middle/Late Pleistocene. The large-scale 100 ka (eccentricity-driven) climate cycles that led to the formation of

Please cite this article in press as: J.R. Lee, et al., The Neogene and Quaternary of England: landscape evolution, tectonics, climate change and their expression in the geological record, Proc. Geol. Assoc. (2017), https://doi.org/10.1016/j.pgeola.2017.10.003

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continental scale ice sheets in North America and Northern Europe resulted in falls of eustatic sea-level by around 120–130 m and appear to have become prevalent from c.1 Ma onwards. Large-scale glacial cycles, occurring on a cyclicity of 100 ka (eccentricitydriven) and being characterised by the development of large ice sheets in northern Europe, including Britain, have only exerted a significant influence on landscape evolution over the past c.1.2 Ma (Sejrup et al., 2005; Lee et al., 2011). Overprinted on these long-term climate cycles and mediumterm glacial/interglacial cycles are abrupt millennial/centennial

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scale climatic events associated with major volcanic eruptions and solar variability. It is often assumed (but not proved) that in Britain and the Northeastern Atlantic, such climate events are triggered by major changes in oceanic circulation such as the North Atlantic Current (i.e. on or off) and the strength of Atlantic Meridional Overturning Circulation (AMOC). Together these exert a strong control on the amount of heat and moisture that the British Isles receives. The main mechanism for disruption or re-organisation of ocean circulation is the dynamic collapse of ice sheets and release of large pulses of freshwater or icebergs into the North Atlantic.

Fig. 2. (a) Alpine folding of Jurassic strata (Purbeck Group) forming part of the Weymouth-Purbeck anticline at Lulworth Stair Hole, Dorset. Bases of key stratigraphic markers beds (Peveril Point Member (P), Ridgeway Member (R), Stair Hole Member (S) and Worbarrow Tout Member (W) (Photo: P937029); (b) Fossil wood (Sequoia) from the Miocene-age Brassington Formation, Kenslow Top Pit, Derbyshire (Photo: P913866).

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Fig. 3. Palaeogeographic evolution of England during the Cenozoic: (a) Mid Miocene showing the main areas of inversion and basinal sedimentation (modified from Knox et al., 2010; Gibbard and Lewin, 2016); (b) Early Pliocene with speculated drainage patterns (modified from Gibbard and Lewin, 2016); (c) Late Pliocene to Early Pleistocene with approximate position of the coastline and reconstructed (solid line) and inferred (dashed line) drainage (modified from Rose et al., 2001); (d) early Middle Pleistocene geography immediately prior to the Anglian glaciation showing the approximate position of the coastline and drainage (modified from Rose et al., 2001; Bridgland et al., 2015); (e) post-Anglian Late Middle Pleistocene (c. MIS 11) palaeogeography showing the Anglian ice limit (red dashed line), reconstructed (blue solid line) and inferred drainage (blue dashed line); (f) Late Glacial palaeogeography showing the maximum extent of Late Devensian ice (dashed red line) and reconstructed (solid blue line) and inferred (dashed blue line) drainage (modified from Gaffney et al., 2007; Bridgland, 2010). Abbreviations: CBB – Cardigan Bay Basin, CB – Cleveland Basin, SGCB – St George’s Channel Basin, SPB Sole Pit Basin, WeB – Wessex Basin, WLB – Weald Basin.

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Consequently, the clearest expression of these abrupt events in environmental records is within glacial stages or during the termination of glacial stages. The effect of abrupt events can be clearly seen in the Greenland ice core records where a large number of sub-stage (interstadial/stadial) cycles have been shown to occur across the last glacial. In Britain and Europe, the clearest expression of landscape/ecosystem response to these abrupt events is during the ‘Lateglacial period’ at the end of the Last Glaciation (Termination 1). 3. Evidence for Neogene and Quaternary landscape change 3.1. The Neogene Period The Neogene Period spans the interval during the Cenozoic Era from the end of the Palaeogene (23.03 Ma) to the beginning of the Quaternary period (2.588 Ma). The Neogene has in-turn been subdivided into two epochs, the Miocene (23.03–5.333 Ma) and Pliocene (5.333–2.588 Ma) and their significance in England is outlined below (Fig. 1a). 3.1.1. Miocene During the Miocene, the primary axis of neotectonic uplift lay parallel to the Alpine Orogenic front which at the time extended from the Massif Central in central France into northern Germany (Bohemian Massif) (Ziegler and Dèzes, 2007; Knox et al., 2010). However, England situated some distance to the north, was also impacted. In southern England and northern France, uplift and folding along pre-existing structures formed the prominent Weald-Artois and Weymouth-Purbeck anticlines and adjacent Hampshire and London Basin synclines (Fig. 2a; Mansy et al., 2003; Ziegler and Dèzes, 2007; Newell and Evans, 2011). Uplift along the Weald-Artois Anticline continued into the Mid-Miocene resulting in localised river incision (Jones, 1999b). Uplift along the axis also led to the initial closure of the Strait of Dover (van Vliet-Lanoë et al., 2010) although other evidence suggests that the ‘Channel Seaway’ which linked the North Sea to the Atlantic remained intermittently open until the Early Pliocene (Gibbard and Lewin, 2003, 2016). Direct geological evidence for the Miocene in England is limited because much of England was undergoing active exhumation (Fig. 3a). This has led to the development of a widespread hiatus (unconformity) extending across Britain (Walsh, 1999), the UK Atlantic margin (Stoker, 2002) and through the North Sea Basin into Fennoscandia (Clausen et al., 1999; Huuse and Clausen, 2001; Hansen and Rasmussen, 2008; Knox et al., 2010). Alluvial-colluvial deposits of Miocene age occur at St Agnes in western Cornwall (Walsh et al., 1987), whilst discontinuous remnants of alluvialfluvial-lacustrine sediments (Brassington Formation) occur as outliers within the Peak District (Pound and Riding, 2016). Palynological assemblages from units within the Brassington Formation indicate a temporal change from pollen produced by sub-tropical, seasonally wet conifer-dominated forest (Fig. 2b) to sub-tropical mixed forest (Pound and Riding, 2016). This has been interpreted as providing evidence for a general climatic cooling which is in-line with climatic evidence for the Miocene from continental Europe (Pound et al., 2012; Pound and Riding, 2016). Elsewhere, marine sands have been recorded in southern England (Balson, 1990) that contain reworked Miocene fossil shark teeth (Carcharocles megladon spp.) (Mathers and Hamblin, 2015). Within and adjacent to the North Sea Basin, the Miocene story is highlycomplex. Localised basin inversion occurred along the Sole Pit Trough (extending northwards into the Cleveland Basin) and also within parts of the Danish and Dutch North Sea sectors (Van Hoorn, 1987; Badley et al., 1989; Corfield et al., 1996; Husse et al., 2001; Rasmussen, 2009). Inversion is likely to reflect the reactivation of pre-existing tectonic structures (e.g. Caledonian and/or Variscan)

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under a north-directed Alpine compressive stress regime (Badley et al., 1989; Mansy et al., 2003) resulting in widespread exhumation and erosion of Mesozoic cover (Green, 1986; Bray et al., 2002). At the same time, enhanced subsidence within the Central Graben combined with marginal uplift and tilting in both Britain and Fennoscandia led to an increase in accommodation space generated within the North Sea Basin (Jordt et al., 1995; Cloetingh et al., 2005). This sediment-infill thickens eastwards away from the British sector (Husse et al., 2001; Overeem et al., 2001; Rundberg and Eidvin, 2005) reflecting higher sediment input from rivers draining the Baltic and Fennoscandia (Eridanos River), Lower Rhine Graben (Proto Rhine), Ardennes, Paris Basin (Proto Somme) and northern Germany and Poland (Overeem et al., 2001; van VlietLanoë et al., 2002; Rasmussen, 2004; Vandenberghe et al., 2004; Preusser, 2008). In Belgium and northeast France, these sediments form distinct coast-parallel sand and gravel ridges (Vandenberghe and Hardebol, 1998). 3.1.2. Pliocene Throughout the Pliocene (5.333–2.588 Ma), the English landscape continued to be regulated by both tectonics (albeit to a lesser extent) and climate. Following an early sea-level rise (Fig. 3b), global sea-levels fell progressively during the Late Pliocene as ice volume in polar areas increased (Fig. 3c). However, uplift and subsidence were arguably the key drivers of long-term landscape development during this interval (Westaway et al., 2002, 2006; Westaway, 2010). A key tipping point occurred at c.3.1 Ma, which according to Westaway (2008), corresponds to a phase of accelerated uplift driven by climatically enhanced weathering and erosion rates. Geological evidence for the Pliocene in England is somewhat limited and largely (but not exclusively) restricted to basinal areas (e.g. the Crag Basin) and in regions situated beyond the limits of glaciation (Jones, 1999a; Rose, 2009; Westaway, 2010). In northern England, Westaway (2009a) has suggested that most of the modern topography was developed following postMiddle Pliocene exhumation with, in general, uplift rates much higher than southern England. According to Westaway (2009a) the principal reason for enhanced uplift in northern England was the greater mobility of the lower crust because of its younger age and buoyancy effects of Palaeozoic igneous intrusions. Erosion of Mesozoic cover-rocks across much of northern England, including the Lake District and the Northern Pennines, occurred during Cenozoic exhumation. The total amount of exhumation is estimated to have totalled between 1,200 and 1,750 m (Holliday, 1993) although local rates may vary (Green et al., 2012). Where this eroded sediment is now stored is an intriguing question (Westaway, 2009a). The Irish Sea is an obvious sediment sink situated adjacent to these exhumed areas. However, unlike during the later Quaternary when the ‘basin’ became a major depositional centre, there are no equivocal Pliocene-age sediments within the ‘basin’ (Jackson et al., 1995). Instead, the basin underwent exhumation and inversion which ended during the Palaeogene (Tucker and Arter, 1987; Williams et al., 2005). Thus during the Pliocene, the Irish Sea was probably part of the UK landmass (Knox et al., 2010) with speculation that the eroded material was transferred to parts of the northern North Sea or the north Atlantic (Westaway, 2009a). Beyond the limits of Pleistocene glaciation in southern England, neotectonic uplift was driven by ongoing northwards-directed Alpine compression. In Devon and Cornwall, several terrestrial or coastal (e.g. rock platforms and raised beaches) erosion surfaces provide evidence for uplift during and after the Pliocene (Robson, 1944; Walsh et al., 1987; Westaway, 2010). Recent work has suggested that following the Mid-Pliocene (but pre late Early Pleistocene) uplift increased progressively between west Cornwall (c.130 m) and the Hampshire Basin (c.150 m) (Westaway, 2010). No

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Pliocene sediments are known within or bordering the English Channel in England (Hamblin et al., 1992) although fluvial and tidal sediments counterparts have been described in northern France (van Vliet-Lanoë et al., 2010). The major basinal area adjacent to England was the Southern North Sea. By the Early Pliocene it is widely believed that the Southern North Sea was separated from the ‘Channel Seaway’ and its link to the north Atlantic was via a northern route situated between Shetland and Norway (Knox et al., 2010; Gibbard and Lewin, 2016). Sedimentation within the basin increased 10-fold during the Pliocene (Kooi et al., 1991; van Wees and Cloetingh, 1996) with an estimated 800 m of sediment deposited within the Dutch part of the Central Graben (Kuhlmann et al., 2004). The generation of accommodation space was driven principally by the reactivation of extensional faulting within the Central Graben (Kooi et al., 1991; Gölke and Coblentz, 1996) combined with

accelerated marginal uplift and tilting including the emergence, during the Late Pliocene, of the Norwegian Shelf (Japsen et al., 2007). Much of the sediment was derived from several large rivers that drained land-masses situated along the eastern margins of the basin. This resulted in the westward progradation of a large wavedominated delta (Overeem et al., 2001; Boenigk and Frechen, 2006; Kuhlmann et al., 2006a; Kuhlmann et al., 2006b). The limited eastwards extent of delta progradation suggests that only minor sediment input was received from rivers draining eastern England (Cameron et al., 1992; Cameron et al., 1993; Rose et al., 2001; Riches, 2012). By comparison with rivers from the continent, rivers in central and eastern England were much smaller, sediment budgets were lower and landscapes relatively stable (Rose, 2010). In eastern England, the westernmost extension of the Pliocene North Sea Basin corresponds to the Crag Basin in East Anglia (Mathers and Hamblin, 2015; Fig. 3c). Sedimentation within the

Fig. 4. Schematic summary of the Mid-Pliocene through to Holocene stratigraphy of eastern England. Top figures show the stratigraphic relationship through the Crag basin (Mid-Pliocene to Early Pleistocene), middle figures show the stratigraphic relationship through the terrace deposits of the Thames and the East Anglia coastal sequences. Bottom show the sedimentary record of the Last Glacial period through to the early Holocene.

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Fig. 5. (a) Non-shelly cross-bedded (lower) and horizontally-bedded sands of the Red Crag Formation, Broom Pit, Suffolk (Photo: P211460). (b) Laminated muds (base), overlain by stratified sands (intertidal) and stratified sands and gravels (beach; ‘Westleton Beds’) comprising the Norwich Crag Formation at Easton Bavents, Suffolk (Photo: P942091). (c) Laminated muds overlain by stratified sands (tidal flats) of the Wroxham Crag Formation, truncated by an ice wedge cast (Barham Arctic Soil), Mundesley, Norfolk. (d) Deformed tidal rhythmites (sand and mud) overlain by shelly freshwater muds (the Sidestrand Unio Bed, Cromer Forest-bed Formation) and marine sands (Wroxham Crag Formation), Sidestrand, Norfolk (Photo: I Candy). (e) Stratified fluvial sands and gravels (Kesgrave Catchments Sub-Group) of the proto-Thames with a cryoturbated rubified palaeosol (Valley Farm Soil) capping the sequence, Stebbing, Essex (Photo: P212609). (f) A channel infilled by freshwater organic muds (Cromer Forestbed Formation) outcropping on the foreshore at Happisburgh, Norfolk (Photo: J Rose).

Crag Basin is recorded by the Crag Group, which although highlyfragmented with numerous hiatuses and unconformities, provides good evidence for the geography of the region from the Early Pliocene through to the late Middle Pleistocene. ‘Crag’ as a term is

derived from the local nomenclature of eastern England and refers to shelly sands and is characteristic of most of the deposits of this basin, reflecting their sub-tidal to littoral marine origin, dominated by relatively high-energy tidal currents and wave processes. The

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earliest preserved sediment infill includes Early-Late Pliocene calcarenites (Coralline Crag Formation) and Late Pliocene, to earliest Pleistocene shelly sands and muds (Red Crag Formation) (Figs. 4 and 5a). These sediments collectively record relatively open-water marine conditions albeit within the context of a progressive shallowing basin (Dixon, 1979; Mathers and Zalasiewicz, 1988; Zalasiewicz et al., 1988; Balson et al., 1993; Hamblin et al., 1997). This shallowing is considered to reflect reductions in accommodation space due to increased sedimentation rates and a fall in global sea-level (Mathers and Zalasiewicz, 1988; Funnell, 1996). Furthermore, the fossil assemblages of the Crag deposits indicate a progressive cooling of the climate and the onset of regular climate cycles. Equivalents to these sediments have been recognised offshore (Cameron et al., 1992) although their onshoreoffshore correlation remains ambiguous (Riches, 2012). The precise origin of the Crag Basin also remains speculative with tectonic (Bristow, 1983; Hopson and Bridge, 1987; Hamblin et al., 1997; Moorlock et al., 2002), erosional (Funnel, 1972; West, 1972; Gibbard et al., 1998) and polygenetic (Riches, 2012) models variably employed to explain its development. Climate has also played a significant role in shaping the landscape of England during the Pliocene by controlling weathering rates and in-turn sediment availability and catchment dynamics (Rose, 2010). Limited eastwards delta progradation in the Southern North Sea (Cameron et al., 1992; Cameron et al., 1993) suggests that landscapes in England were relatively stable, with sediment budgets and catchment dynamics low (Rose et al., 2001). In general terms, Pliocene climates in northwest Europe are considered to have been a little warmer than present day and this signal is replicated by data from the Crag Basin (Haywood et al., 2000; Johnson et al., 2000; Williams et al., 2009). However, in a review of palaeoclimate analyses of different fossil groups, Williams et al. (2009) found conflicting evidence with some groups indicating warmer, and others suggesting cooler climatic conditions. 3.2. The Quaternary Period The Quaternary Period is the most recent of three geological periods during the Cenozoic Era and spans the past 2.588 Ma of geological time (Gibbard et al., 2010a). It is sub-divided into the Pleistocene (2.588 Ma–0.0117 Ma) and Holocene (0.0117 Ma – present day) epochs. 3.2.1. Early Pleistocene The Early Pleistocene corresponds to the time interval between 2.588 Ma and 0.78 Ma. Uplift and subsidence played an important long-term role in driving landscape development during the Early Pleistocene by regulating accommodation space that enabled deposits of this age to be deposited and preserved (i.e. the Crag Group of East Anglia). Westaway et al. (2002) and Westaway (2009b) have identified an acceleration of uplift rates in eastern and central England at c.2.0 Ma and argue that this corresponds to a climatically-triggered acceleration in erosion rates. Beyond the Crag Basin, very few sites record significant Early Pleistocene deposition implying that terrestrial conditions prevailed across much of England. The time-interval of the Late Pliocene through the Early Pleistocene is, in the benthic d18O signal, characterised by an intensification of the magnitude and frequency of glacial/ interglacial cycles (Lisiecki and Raymo, 2007). Whilst the Late Pliocene was characterised by subdued, low magnitude 20 ka climate cycles, the Early Pleistocene saw the onset of more pronounced 40 ka cycles, and it is evidence for this increase in the magnitude of climate cooling that characterises much of the sedimentary record of this interval in Britain (Fig. 7a). The later

deposits of the Red Crag Formation and succeeding Norwich Crag Formation are proposed, based on biostratigraphy and palaeomagentic dating, to span the Late Pliocene and Early Pleistocene interval through to c.1.8 Ma (Fig. 4; Mathers and Zalasiewicz, 1988; Mathers and Hamblin, 2015). The Norwich Crag Formation in particular contains sedimentary evidence for cyclic changes in sea-level with sub-tidal, intertidal and littoral facies (Fig. 5b) present (Zalasiewicz and Mathers, 1985; Zalasiewicz and Gibbard, 1988; Hamblin et al., 1997; Richards et al., 1999). Pollen analysis of these deposits also record fluctuations between deciduous woodland and open heath in adjacent onshore environments (Funnell and West, 1962; Norton and Beck, 1972; Funnell et al., 1979). In neither the sedimentary nor palaeoecological records is there any suggestion that arctic tundra environments or widespread permafrost occurred during the earliest Pleistocene. Gravel clast lithology of the late Red Crag and Norwich Crag formations are dominated by flint (>95%) derivedlocally from Cretaceous or reworked from Palaeogene strata (Rose et al., 2001, 2002). They indicate that the rivers that fed the Crag Basin at this time were either local systems with spatiallyrestricted catchments or, more probably, larger systems that did not possess sufficient stream power to transport coarse-grained bed-load from the headwaters to the lower reaches of their catchments (Rose et al., 2001; Rose, 2009). As the fine-grained facies of the Norwich Crag Formation contain palynomorphs derived from strata in the Welsh Borders, English Midlands and the Pennines, the latter scenario is most probable (Riding et al., 1997, 2000; Moorlock et al., 2002). Norwich Crag Formation deposits do, however, contain far-travelled heavy minerals within intertidal clays (e.g. at Easton Bavents) that indicate that icebergs may have been present in the Southern North Sea at the time (Funnell and West, 1962). This interpretation is supported by records of iceberg ploughmarks within 3D seismic data from various sectors of the North Sea Basin (Graham, 2007; Kuhlmann and Wong, 2008; Dowdeswell and Ottesen, 2013). Traditionally held views that glaciation was not a significant feature of the Early Pleistocene landscape appear to be out-dated. A growing body of evidence now argues that glaciers may have been active in highland areas (and adjacent basins) of western and northern Britain, Ireland and Fennoscandia periodically throughout the Early Pleistocene (Stoker et al., 1994; Sejrup et al., 2005; Ottesen et al., 2009; Lee et al., 2011, 2012; Thierens et al., 2012). Significantly, during the earlier part of the Early Pleistocene the Crag Basin was fed by several major river systems, including the proto-Thames, that drained eastwards across central and eastern England (Fig. 3c). However, the scale and extent of the protoThames during the Early Pleistocene has proven controversial. Originally, the course of the proto-Thames was traced northeastwards across East Anglia into north Norfolk with, what is now called the Bytham River, believed to have been a major left-bank tributary (Hey and Brenchley, 1977; Hey, 1976, 1982, 1996; Green et al., 1980; Green and McGregor, 1990, 1999; Rose et al., 1996a, 1996b; Ashton et al., 2011). This model has been questioned with, as an alternative model, the proto-Thames deposits of Norfolk reinterpreted as marine sediments (part of the Wroxham Crag Formation) (Hamblin and Moorlock, 1995; Hamblin et al., 1996; Briant et al., 1999; Rose et al., 2001, 2002; Lee, 2009) and the Bytham and proto-Thames rivers draining separately into the North Sea Basin (Hamblin and Moorlock, 1995; Riding et al., 1997; Rose et al., 2001, 2002; Moorlock et al., 2002; Lee et al., 2006; Rose, 2009). Further debate also surrounds the origin of the Nettlebed Formation. Conventionally, this deposit has been interpreted as an early unit of the proto-Thames (Gibbard, 1985; Whiteman and Rose, 1992) and correlated with both the Red and Norwich Crag formations (Crag Group) due to the dominance of flint gravel clasts (Rose et al., 2001). The interpretation of the Nettlebed Formation

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as an early proto-Thames deposit has recently by challenged by Worsley (2016) who has instead interpreted it as a slope deposit. Many have argued that a regional-scale sedimentary hiatus developed within the Crag Basin between c.1.8–0.78 Ma (West, 1977; West, 1980; Jones and Keen, 1993; Kemp et al., 1993). This view was based upon the absence of biostratigraphic assemblages and indicator species that characterise this interval in continental Europe but have yet to be identified within the British record. The origins of this hiatus remain speculative but are likely due to a combination of accelerated marginal uplift (Westaway et al., 2002) and a dramatic reduction in basin accommodation space. Evidence for this hiatus include the development of palaeosols, such as the Valley Farm (temperate) and Barham Arctic (periglacial) soils (Fig. 5b, c and e), which are superimposed upon both Crag Group deposits and river terrace surfaces across East Anglia (Rose and Allen, 1977; Rose et al., 1985a; Kemp et al., 1993; Rose, 2015). Not only do these soils provide evidence for long-term landscape stability, but detailed soil micromorphological analyses reveals the presence of composite periglacial-temperate soil profiles that developed over multiple glacial-interglacial cycles (Kemp, 1987; Whiteman and Kemp, 1990; Kemp et al., 1993; Read et al., 2000). Other geological evidence for this time-interval includes older fluvial beds (Siliceous Member) within the cave succession at Westbury-sub-Mendip (Bishop, 1975, 1982; Andrews et al., 1999) and the terrace archive of the proto-Thames (Gibbard, 1985; Whiteman and Rose, 1992; Bridgland, 1994, 2006, 2010; Westaway et al., 2002). The absolute age of individual terraces within the proto-Thames remain unclear because these deposits are relatively poor in fossil material and contain little material suitable for dating (Whiteman and Rose, 1992; Rose et al., 1999; Westaway et al., 2002). However, the lithology of these deposits records how the proto-Thames catchment has evolved through time. The switch from locally-dominated (>95%) lithologies (Nettlebed Formation; cf. Worsley, 2016) to far-travelled lithologies (upto 50%; Sudbury Formation) records a major Early Pleistocene expansion of the

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proto-Thames catchment with headwater connections to both North and South Wales, the English Midlands and the Pennines (Green et al., 1980). A later (c.0.8 Ma) return to locally-derived lithologies (Colchester Formation), with a persistent far-travelled component (c.30-5%), suggests a dramatic reduction in catchment size possibly as a result of an early phase of glaciation or river capture (Whiteman and Rose, 1992, 1997; Bridgland, 1994; Rose et al., 1999, Rose, 2009). Recent research suggests that sedimentation resumed within parts of the Crag Basin slightly earlier than previously considered (c.1.0–0.9 Ma). This was probably because of the increased availability of accommodation space driven by a second phase of accelerated marginal uplift and basin subsidence (Rose et al., 2002; Westaway et al., 2002; Westaway, 2009b, 2017). Coastal and shallow marine deposits encompass older units of the Wroxham Crag Formation and to-date, have been recognised at and beneath modern beach-level along the north Norfolk coast between West Runton and Happisburgh (West, 1980; Briant et al., 1999; Allen and Keen, 2000; Larkin et al., 2011). Happisburgh is particularly significant because sediments, deposited under a boreal climate, are believed to contain Mode 1 archaeology (core and flake technology) which may represent the oldest (c.0.99–0.8 Ma) known evidence for human occupation north of the Alps (Fig. 6; Parfitt et al., 2010). However, some caution is required in interpreting the archaeology from Happisburgh (Bridgland and White, 2014). Firstly, the limited size of the archaeological assemblage (78 artefacts) cannot unequivocally validate its Mode 1 interpretation (Bridgland and White, 2014). Secondly, whilst the Happisburgh sequence may be as old as MIS 25-21 (Parfitt et al., 2010; Preece and Parfitt, 2012), elsewhere it has been interpreted as being much younger (MIS 15c; Westaway, 2011). Some have argued the presence of far-travelled vein quartz within these deposits (Parfitt et al., 2010) to indicate correlation with protoThames deposits (Sudbury Formation). However, the ‘Happisburgh-on-Thames’ concept isn’t universally accepted, partly

Fig. 6. Comparison between key phases of Palaeolithic archaeological occupation of Britain and the LR04 benthic record (see Ashton et al., 2011; Pettitt and White, 2012 for details). Different chronologies have been suggested for the earliest (pre-Anglian or MIS 12) archaeology with researchers such as Parfitt et al. (2005, 2010 and Hosfield (2011) suggesting a long chronology (spanning the Early to early Middle Pleistocene). In this model it is likely that most, if not all, pre-Anglian Acheulian sites can be correlated with MIS 13 but there is a possibility that some may date to late MIS 15 (see Candy et al., 2015a, 2015b for discussion). In contrast Westaway (2010, 2011) has proposed that all preAnglian archaeological sites can be accommodated into MIS 13 and 15. MIS 7 contains multiple substages that all appear to be of fully interglacial conditions (Candy and Schreve, 2007). However, most archaeological sites would appear to date to the later part of this isotopic stage although at West Thurrock it is likely that the Levallois artefacts correlate to the very end of MIS 8. Levallois sites routinely contain hand axes but here these are not ascribed to either Acheulian or Mousterian types.

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Fig. 7. Patterns of Quaternary climate forcing observable in various long climate archives. (a) The Lisiecki and Raymo (2005) stacked benthic d18O record for the past 3 million years and the progressive evolution of the magnitude and frequency of climate cycles. (b) Graph showing the EPICA Dome C temperature anomaly record (Jouzel et al., 2007) of the past 800,000 yrs and the step like shift in the intensity of glacial/interglacial cycles that occurs from MIS 12 (450,000 yrs) onwards. (c) The Greenland (NGRIP, 2004) record of the past 125,000 yrs, highlighting the abrupt, high magnitude climatic events (Dansgaard-Oeschger, or D/O, cycles) that occur during this interval. (d) The Last Glacial to Interglacial Transition (LGIT) as expressed in NGRIP and the millennial/centennial scale climatic events that are seen in this period.

because the vein quartz could be derived from other fluvial sources (Westaway, 2011). At several other localities in north Norfolk, discrete horizons of far-travelled heavy minerals and erratic clasts from northern Britain and Fennoscandia have also been recorded within these coastal deposits. They are thought to have melted-out from grounded icebergs indicating the presence of glaciers

bordering and within the North Sea Basin (Lee, 2003; Larkin et al., 2011). 3.2.2. Middle Pleistocene The benthic d18O signal records the occurrence of two major climatic transitions that are significant for the climate of the

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Middle Pleistocene (0.78–0.13 Ma). The first of these is the Middle Pleistocene Transition (MPT), which reflects the increasing dominance of the more extreme 100 ka glacial cycles (Fig. 7a; Clark et al., 2006; Lawrence et al., 2010; McClymont et al., 2013; Elderfield et al., 2012). In some records, this appears as a sudden shift from 40 ka to 100 ka cycles at c. 1 Ma (Elderfield et al., 2012), whilst in others it appears to be a more gradual and progressive transition between c.1.2–0.6 Ma (McClymont et al., 2013). Regardless of the precise expression of the MPT, it is only after this event that high-magnitude and long-duration glacial stages occur. With respect to the British record, there is no evidence for continentalscale glaciation or widespread periglaciation until after the MPT (Lee et al., 2012). The second transition, the mid-Brunhes Event (MBE), represents the shift between MIS 13/15 from relatively subdued glacial/interglacial 100 ka cycles to climate cycles characterised by a greater intensity in both the peak of interglacial warming and the extreme of glacial cooling (Fig. 7b; Jansen et al., 1986; EPICA, 2004; Lang and Wolff, 2011). Evidence for the MBE in the British Isles is less convincing as many of the warmest interglacial temperature reconstructions in this region are recorded in interglacial deposits of the early Middle Pleistocene (0.78–0.45 Ma), whilst deposits of the late Middle Pleistocene (0.45–0.13 Ma) contain evidence for widespread periglaciation ( Candy et al., 2010, 2011; Candy and McClymont, 2013). Prior to the onset of lowland glaciation within England, the Crag Basin continued to be the primary depositional centre with shallow marine and coastal deposits corresponding to younger units of the Wroxham Crag Formation (Rose et al., 1996a, 1996b; Briant et al., 1999; Pawley et al., 2004; Lee, 2009; Fig. 5c and d). They intercalate with river terrace sequences of the proto-Thames (Colchester Formation; Fig. 5e) and Bytham, plus floodplain deposits (Cromer Forest-bed Formation; Fig. 5f) associated with several smaller rivers in north Norfolk that drained into the Southern North Sea (Figs. 3d and 4; Whiteman and Rose, 1992; Lewis, 1993; Preece and Parfitt, 2000; Rose et al., 2001, 2002; Lee et al., 2006; Rose, 2009). Clast lithologies within the proto-Thames and Bytham deposits, plus their coastal Crag Group equivalents, contain a mixture of durable local and far-travelled types (e.g. flint, vein quartz, quartzite and chert). Clast populations reflect catchments that were subjected to active weathering and slope processes, and rivers that were both increasingly effective at recycling coarse material (Whiteman and Rose, 1992; Rose et al., 1999, 2001; Lee et al., 2008; Rose, 2009) and removing (by fluvial abrasion) the non-durable gravel component (Rose et al., 1976; Rose and Allen, 1977). Interglacial fluvial deposits by contrast, record low-energy deposition in the lower reaches of large river catchments, primarily the Bytham, and more local rivers (West, 1980; Preece and Parfitt, 2000; Candy et al., 2006; Lee et al., 2006; Gibbard et al., 2010b). The sedimentology of these deposits is often complicated by the lag that occurs between temperature amelioration and sea-level rise during glacial terminations. Consequently, most of these sites record the early part of the interglacial as fully-continental fluvial sedimentation whilst the later part of the interglacial is recorded by tidal/shallow marine sedimentation, reflecting the marine transgression that floods the lower part of the river system as high stands are achieved (Lee et al., 2006). Key drivers of landscape change during this period were the frequent and marked changes in climate that drove major changes in sea-level, palaeogeography and in-turn surface geological processes (Allen and Keen, 2000; Lee et al., 2006). Climatic indicators including sedimentology, palaeosols, isotopes and palaeontology suggest marked extremes in climate between Mediterranean-style temperate climates (Kemp et al., 1993; Candy et al., 2006) and periglacial cold climates (Rose et al., 1985a; Kemp et al., 1993; Fish et al., 1998; Lee et al., 2003, 2006; Larkin et al.,

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2014). Previously, this preglacial interval was defined as a single ‘Cromerian’ interglacial (West, 1980). However, as many as five different interglacial events (the ‘Cromerian Complex’) have been identified based on biostratigraphy (Preece and Parfitt, 2000, 2012; Preece, 2001; Stuart and Lister, 2001) and amino acid racemisation (AAR) (Penkman et al., 2011). One of the most significant consequences of Cenozoic exhumation and climate change was that it degraded the landscape (and properties of the shallow sub-surface) for exploitation during periods of glaciation (e.g. Phillips et al., 2013). The most extensive of these glaciations corresponds to the Anglian Glaciation (Fig. 3e; Perrin et al., 1979; Bowen et al., 1986; Clark et al., 2004; Lee et al., 2011). Deposits occur extensively across central and eastern England (Fig. 4; Bowen et al., 1986; Hart, 1990; Ehlers and Gibbard, 1991; Lee, 2001; Phillips et al., 2008; Rose, 2009; Lee et al., 2013, 2017) with its age assigned to MIS 12 (480–430 ka) using geochronology and biostratigraphy (Candy et al., 2014). Conventionally, it has been correlated with the Elsterian Glaciation of continental Europe (Zagwijn, 1986; Ehlers and Gibbard, 1991). Historically, there has been considerable debate surrounding the number of Middle Pleistocene glaciations that have affected England. By their nature, glaciations are highlyerosive, resulting in sequences that frequently preserve only evidence for the most extensive (i.e. MIS 12) and recent glaciation (i.e. MIS 2). Evidence has been presented for a pre-Anglian glaciation in East Anglia (c. MIS 16) (Hamblin et al., 2000; Lee et al., 2004; cf. Preece et al., 2009) and various post-Anglian/preDevensian glaciations in the Midlands and Eastern England (West and Donner, 1956; Catt and Penny, 1966; Straw, 1979, 1984; Sumbler, 1995; Hamblin et al., 2005; White et al., 2010, 2017a; Davies et al., 2012a; Langford, 2012; Bridgland et al., 2014, 2015; Powell et al., 2016). The Anglian glaciation had several major impacts on both the landscape and geological record of Britain. Glaciation resulted in the marked erosion of the pre-existing bedrock landscape (Gallois et al., 1999; Clayton, 2000) with eroded materials redistributed laterally into more marginal glacial areas (Fig. 8; Perrin et al., 1979; Rose, 2009; Scheib et al., 2011; Lee et al., 2017). In contrast to earlier fluvial and shallow marine Pleistocene deposits, a distinctive feature of the lithology of glacially-sourced materials is the marked increase in low-durability gravel clasts (e.g. Rose et al., 1976; Rose and Allen, 1977). This reduction in compositional ‘maturity’ reflects the increased preservation of low-durability gravel lithologies due to increased sediment budgets, shorter transit distances and more rapid sedimentation rates. Leeder (2008) has speculated that glacial excavation of the Fen Basin and lowering of the chalk escarpment has led to uplift along the eastern Fen margin influencing post-glacial drainage development. The post-glacial drainage pattern of central and eastern England is arguably one of the most obvious relics of the Anglian Glaciation (Bridgland, 2010). Preglacial drainage in central and eastern England was typically either over-ridden or diverted by ice (Gibbard, 1977, 1985; Bridgland, 1994, 2010; Rose, 1994). Reestablishment of drainage following deglaciation resulted in the formation of several smaller river catchments including those of the modern Thames and Trent and rivers that drain into and through the Fen Basin (Rose et al., 1985b; Bridgland, 1994, 2010; White et al., 2010; Banks et al., 2012; Bridgland et al., 2014; Westaway et al., 2015). This fundamental drainage change has resulted in the Crag Basin ceasing to exist as a major sediment sink causing regional basin subsidence to stop (Lee, 2015). The Anglian Glaciation also records the first time since the Late Miocene that Britain became physically-separated from continental Europe (Preece, 1995). Prior to the Anglian, Britain and continental Europe were joined by the Weald-Artois anticline. However, catastrophic drainage of an extensive ice-dammed lake basin, which formed in

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Fig. 8. Patterns of Middle Pleistocene glaciation (‘Anglian’ and possible ‘pre-Anglian’) in East Anglia and the Midlands showing the main zones of substrate erosion and accretion (modified from Lee et al., 2017). This new model of glaciation for the region employs a tectonostratigraphic approach for classifying the glacial evolution. Cartoons ‘c-f’ record four main stages of ice-marginal advance and retreat attributed to the main ‘Anglian Glaciation’. Cartoons ‘a-b’ record two earlier ice-advances into northeast Norfolk, which are separated from the ‘Anglian Glaciation’ by a regionally extensive unconformity with discrete evidence for intervening phases of widespread landscape erosion and stabilisation, sea-level change and soil development (Lee et al., 2017). The precise timing of this earlier glaciation, the so-called ‘Happisburgh Glaciation’ remains the matter of much discussion and debate (e.g. Lee et al., 2004; Hamblin et al., 2005; Rose, 2009; Preece et al., 2009; Westaway, 2009b).

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the Southern North Sea, led to meltwater incision through the structural high (Smith, 1985; Gibbard, 1988, 1995). Formation of the Strait of Dover was significant for two reasons. Firstly, it meant that the Southern North Sea was connected to the North Atlantic via the English Channel during periods of high global sea-level (Toucanne et al., 2009). Secondly, that Britain became an island

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during several subsequent interglacial sea-level high stands forming temporary migration barriers to ancient humans and mammals (White and Schreve, 2000). With the exception of the Hoxnian Interglacial (MIS 11), the overall record of post-Anglian interglacial stages in England is strongly-biased towards fluvial sediments, with rivers under

Fig. 9. (a) Photograph Brighton-Norton Raised Beach, Black Rock (Brighton) showing the ancient chalk cliff line, raised beach deposits and overlying periglacial slope deposits; the deposits were formed during the MIS 7-6 transition and form part of the West Sussex Coastal Plain (Photo: P212513). (b) Silt and clay couplets (varves) deposited in a lake basin through the transition from the Anglian Glaciation (MIS 12) into the following Hoxnian Interglacial (MIS 11), Marks Tey, Essex (Photo: P212037). (c) Late Devensian subglacial bedforms (lineations and drumlins) near Appleby, Eden Valley, shown within a NEXTMAP Digital Surface Model (DSM) [NEXTMap Britain elevation data from Intermap Technologies]. A first flow set shows initial eastwards-directed ice-flow towards the Stainmore Gap situated to the east (see Evans et al., 2009 for further information). A second SE-NW orientated flow set record later NW-directed ice flow towards the Solway lowlands (see Evans et al., 2009 for further information). (d) Late Devensian till (Skipsea Till) overlain by outwash sands and gravels, Skipsea, East Yorkshire (Photo: J Lee). (e) The Ringstead Dry Valley, Hunstanton (Norfolk), formed by the catastrophic drainage of an ice-dammed lake at the southern limit of the North Sea Lobe of the last British-Irish Ice Sheet (Photo: P210717).

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temperate climates characterised by low energy single-thread, meandering systems. This results in interglacial sequences accumulating in abandoned channel features that have limited accommodation space. Consequently, very few sites exhibit a complete sedimentary record of an entire interglacial and the majority of Middle Pleistocene interglacial deposits therefore represent short-lived “snapshots in time” (Candy et al., 2014, 2016a). The most complete post-Anglian late Middle Pleistocene record of interglacials are recorded within the terrace sequence of the Lower Thames (Fig. 4; Bridgland, 2006). A combination of morpho-, litho- and bio-stratigraphy, associated with AAR correlation and absolute dating, has led to the suggestion that one terrace unit was/has been formed per glacial/interglacial cycle (Bridgland, 2000; Schreve, 2001; Roe et al., 2009; Penkman et al., 2011). The model of terrace development for the Lower Thames and several other major lowland river systems (e.g. Trent and Solent) includes the implicit assumption that fluvial response to climate change is occurring against a background of long-term tectonic uplift (Bridgland, 2000; Maddy and Bridgland, 2000; Westaway et al., 2002, 2006; Howard et al., 2007; Lane et al., 2008). Uplift rates accelerated following the passing of a major climatic threshold at c.0.9 Ma albeit with higher rates inferred from the early Middle Pleistocene compared to the late Middle Pleistocene (Westaway et al., 2002; Lane et al., 2008; Westaway, 2009b; Farrant et al., 2014). Uplift results in the vertical separation of terrace landforms of different ages, leading to the view that their elevation is a function of the landform’s age. Similar principals of terrace generation have been applied to the development of raised beaches sequences including the West Sussex Coastal Plain (WSCP) (Bates et al., 1997; Briant et al., 2006; Westaway et al., 2006). Raised beaches (Fig. 9a) within the WSCP have been identified that can be correlated to each interglacial from MIS 13 onwards (Bates et al., 1997). In the case of the WSCP, it is suggested that progressive uplift was influenced by vertical slip on the blind reverse fault (Alpine-related) situated beneath the Portsdown Anticline (Westaway et al., 2006). Whilst river terrace sequences are found in association with most lowland river systems and isolated raised beach deposits occur around the coastline, it is the records of the Lower Thames and the WSCP that provide the most detailed and complete record of long-term interglacial/glacial climate forcing in England (Bates et al., 1997; Bridgland, 2000). It is important to note that whilst many long climate records from around the world indicate that interglacials of the early Middle Pleistocene are consistently cooler than those of the Holocene and late Middle and Late Pleistocene (as per the MBE) this does not appear to be true for the British interglacial record (Candy et al., 2010). The key thermophilous fossil species that are routinely used in the British Isles to indicate interglacial climates that were warmer than the present day are found as frequently before the MBE as after (Coope, 2010; Candy et al., 2010). In-fact many of the warmest interglacial climatic reconstructions derived from the British sequence are found within deposits of the ‘Cromerian Complex’ (Candy and McClymont, 2013). Whereas the fluvial record dominates much of the interglacial record, the sedimentary sequences of the Hoxnian interglacial (correlative of the Holsteinian, MIS 11) are an exception to this. These sequences are largely lacustrine sequences found within remnant kettle holes and deep bedrock-incised ‘tunnel valleys’ that formed during and after deglaciation (Turner, 1970; Coxon, 1985; Hart and Peglar, 1990; Boreham and Gibbard, 1995; Rowe et al., 1999; Preece et al., 2009). These lake basins accumulated sediments spanning the late Anglian and part (or in some cases all) of the subsequent Hoxnian interglacial (Fig. 9b; Candy et al., 2014). In contrast interglacials that are recorded primarily in fluvial deposits are characterised by highly fragmented records, with

stages such as the Last Interglacial (the Ipswichian or MIS 5e) containing a rich record of the early part of this warm episode but limited evidence for deposits correlating to the later part of this interglacial (Candy et al., 2016a). The Middle Pleistocene contains a wealth of evidence for Palaeolithic occupation (Fig. 6; Wymer, 1999; Hosfield, 2011; Pettitt and White, 2012; Candy et al., 2015a). In the pre-Anglian (
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(MIS 9) in the lower Thames terrace sequence provide good lithostratigraphic records for these interglacials and highlight well the diversity of archaeology that is found within them (Bridgland, 1994; Schreve et al., 2002; White et al., 2013). In both of these interglacials, there is good evidence for a return to Mode I archaeology, known as the Clactonian industry, in the earliest parts of both of these interglacials (White and Schreve, 2000). The Clactonian is then replaced by Mode II Acheulian industries in the middle parts of both MIS 11 and 9. The later part of both interglacials sees technological innovation with the appearance of twisted ovate hand axes in the later part of MIS 11 and the first appearance of Levallois technology in the later part of MIS 9 (White and Schreve, 2000; Ashton et al., 2011; Bridgland and White, 2014, 2015). In Britain, studies of handaxe morphology have advanced to a level where clear temporal patterns in technology and culture can now be recognised, highlighting human-climate interactions and changes in cultural preferences and practices over the past 500,000 years (Bridgland and White, 2015). Thus, during the Middle Pleistocene, the British geological record supports a general pattern of human colonisation and occupation during interglacials and abandonment during glacials (Ashton et al., 2011). This is true until the Middle/Late Pleistocene boundary where humans appear to be absent from Britain during MIS 6 and do not return until the Middle of the last glacial (MIS 3), implying that humans did not recolonise this region during the warm climates of MIS 5e (Ashton and Lewis, 2002; Lewis et al., 2011). 3.2.3. Late Pleistocene (MIS 5e-2) The nature of sediment and landform preservation potential means that the geological record of the Late Pleistocene (effectively the Last Interglacial/Glacial cycle, 0.13–0.0117 Ma) is more rich and diverse than any proceeding climate cycle (Bell and Walker, 2005). Furthermore, the Late Pleistocene encompasses the optimum time-interval where OSL and U-Series geochronology can be employed (although both can often be effective at time-scales extending back into the Middle Pleistocene) and to a more limited extent radiocarbon dating (back to c.50–40 ka) (e.g. Briant et al., 2004; Higham et al., 2006; Gilmour et al., 2007; Dinnis et al., 2016). The body of literature that describe sites and synthesise data from this time interval in Britain is so extensive that a detailed review is beyond the scope of this paper. Consequently, the Late Pleistocene will here be divided into four distinct time intervals: (1) the Last Interglacial (MIS 5e); (2) the Last Glacial (MIS 5d to 3): (3) the Last Glacial, Dimlington Stadial (MIS 2); and (4) the Lateglacial (the earliest part of MIS 1). Whilst discussing earlier parts of the Quaternary record the climate drivers under consideration were orbital cycles or longer-term transitions. During the Late Pleistocene, it is becoming increasingly clear from ice core records, amongst other archives, that much of the last glacial period was characterised by abrupt, high-magnitude climate events operating on millennial and centennial timescales (Fig. 7c; Dansgaard et al., 1993; NGRIP, 2004). This is particularly true for MIS 3 and the Lateglacial (NGRIP, 2004). Whilst it is likely that such events characterised most glacial stages of Middle Pleistocene, the availability of higher-resolution dating techniques and a richer evidence base makes investigating the expression of such events in the British terrestrial record more feasible. However, despite the abundant evidence for extreme abrupt events in the North Atlantic region during MIS 3, in the British Quaternary record such events, and system response to such events, are only clearly seen during the Lateglacial (Atkinson et al., 1987; Mayle et al., 1999; Candy et al., 2016b). 3.2.3.1. The Last Interglacial (MIS 5e). The Last Interglacial, the Ipswichian (Eemian; MIS 5e), is represented in the British Isles, primarily in fluvial and cave sequences and a number of raised

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beach deposits (Coope, 2001; Currant and Jacobi, 2001, 2011; Candy et al., 2016a). The Last Interglacial sequences on the Gower in neighbouring Wales are of particular interest in that, as well as preserving cave and shoreline deposits in places these are overlain by sediments that provide detailed, albeit fragmented, evidence for the subsequent glacial episode (Campbell and Bowen, 1989; Hiemstra et al., 2009). In this region litho-bio-stratigraphic records are available for much of the Last Interglacial/Glacial cycle. Across much of southern and eastern England it is, however, the fluvial sediment archive, that provides the most detailed evidence for the Ipswichian (e.g. Preece, 1999; Gao et al., 2000; Briant et al., 2006; Candy et al., 2016a). As discussed earlier, the fragmentary nature of fluvial sediments means that no site in Britain is currently known to contain a complete record of the Ipswichian (West, 1957). The majority of sites record the early part of the Ipswichian (pollen zones I and II) and these are typified by a rich assemblage of thermophilous taxa (flora and fauna), including Hippopotamus (Coope, 2001; Currant and Jacobi, 2001; Candy et al., 2010, 2016a). This species is often used as an indicator species for the Ipswichian as this is the only interglacial of the past 0.45 Ma when it appears to have been part of the British fauna (part of the distinctive Joint Mitnor Mammal Assemblage Zone (MAZ) argued to be characteristic of MIS 5e in Britain by Currant and Jacobi, 2001). At a number of cave sites the U-Series dating of speleothems associated with the Hippopotamus fauna yield radiometric ages consistent with the attribution of this assemblage to MIS 5e (Currant and Jacobi, 2001, 2011). Palaeoclimatic reconstructions for the Ipswichian in England (based on associated flora and fauna), routinely provide evidence for climates that were warmer than the present day (Zagwijn, 1986; Coope, 2001, 2010; Candy et al., 2016a). At key sites where particularly rich fossil assemblages are found, such as within river terrace deposits beneath Trafalgar Square in central London, the climatic reconstructions indicate a climate that was warmer than at any point within the current interglacial (Candy et al., 2016a). It is widely acknowledged that the palaeoecological evidence of the Ipswichian suggests that this, along with interglacials of the ‘Cromerian Complex’, corresponds to one of the warmest interglacials known to have occurred in Britain over the past 0.8 Ma (Candy et al., 2010). One significant feature of the Last Interglacial is the apparent absence of human occupation during this interval (Lewis et al., 2011). Whilst human artefacts are known from every interglacial since MIS 13, none can be securely attributed to the Last Interglacial (Fig. 6; see also Wenban-Smith et al., 2010). The reason for human absence in Britain during the Ipswichian is currently unknown. 3.2.3.2. Last Glacial (MIS 5d-3). In the Greenland ice core records the early to middle part of the Last Glacial (MIS 5d-3) is characterised by a complex series of interstadial and stadialscale climatic events, although these are not clearly expressed in the British record. In Britain this interval is recorded within numerous fluvial sequences that are typically characterised by braided river deposits (e.g. Maddy et al., 1998; Briant et al., 2004) and by faunal assemblages derived from cave sediments (Currant and Jacobi, 2001, 2011; Jacobi et al., 2006; Pettitt and White, 2012). Effectively, the post-Ipswichian MIS 5 assemblages (represented by the Bacon Hole Mammal Assemblage Zone (MAZ)) retain temperate characteristics but loose the strongest evidence for peak warmth, such as Hippopotamus. The very final part of MIS 5 (most probably late MIS 5a) are characterised by the development of an open landscape and an assemblage of much lower diversity (the Banwell Bone Cave MAZ) dominated by reindeer, bison, wolverine and a very large form of brown bear (Currant and Jacobi, 2001). The Banwell Bone Cave MAZ was previously attributed to MIS 4, the first major cold event of the last glacial, but re-dating of this

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assemblage by U-series dating has re-assigned it to late MIS 5a (Currant and Jacobi, 2001, 2011; Gilmour et al., 2007). Currently, no known fossil assemblage can be attributed to MIS 4, a fact that may reflect the extreme cold of the time-interval or be a reflection of preservation potential. MIS 3 in Britain is characterised by the Pin Hole MAZ, and exhibits species characteristic of open ground and cooler conditions. These include horse, woolly mammoth, woolly rhinoceros, bison, reindeer, hyena and artefactual evidence for the return of humans into the British Isles (Currant and Jacobi, 2001, 2011; Boismier et al., 2012; Schreve et al., 2012). On the basis of lithic technology and also the dating of human remains, both Neanderthal and anatomically modern Homo sapiens appear to have been present in Britain during MIS 3 (Jacobi and Higham, 2008; Higham et al., 2011). However, the age of the dated Homo sapiens remains indicate a late MIS 3 age whilst OSL dating of Mousterian assemblages at Lynford are suggestive of ages at the very onset of MIS 3 (Boismier et al., 2012). It is apparent that the sedimentary record of Britain contains evidence for numerous warm intervals of interstadial status during the Last Glacial period (Coope et al., 1997; Penkman et al., 2011). The climatic complexity of this period can be highlighted through coleopteran-based temperature reconstructions from a number of sites that fall into the post-Ipswichian but preDimlington interval, although their absolute ages or indeed stratigraphic order cannot currently be confidently reconstructed (Coope et al., 1997). Whilst many of these reconstructions highlight the cool conditions of many of the interstadials, at sites such as Isleworth (London), deposits of the Thames that are suggested to be of MIS 3 age indicate summer temperatures as warm or warmer than the present day (Coope et al., 1997). The chronology of these interstadials is much debated, primarily because they occur at or beyond the limit of radiocarbon dating (Bowen et al., 1989; Penkman et al., 2011). Although such palaeoecological assemblages may provide evidence for numerous interstadial events, consistent with the climatic record preserved in the Greenland ice cores, there is limited evidence within sedimentary sequences to suggest a noticeable geomorphic response to such events (van Huissteden et al., 2001). In multiple catchments in southern and eastern England (i.e. the Upper Thames and the Nene) braided river deposits, with interbedded ice wedge casts, have been OSL dated to MIS 3 but contain no clear evidence for fluvial system response to abrupt events (Maddy et al., 1998; Briant et al., 2004). Whether this is because these events were of insufficient magnitude or duration to cause the fluvial system to cross a threshold or produce substantive change is currently unclear. 3.2.3.3. Last Glacial (Dimlington Stadial/MIS 2). The presence of glaciers in England during earlier stages of the Last Glacial stage has been much debated although it is now widely accepted that the Dimlington Stadial (c.28–18 ka) (Rose, 1985) was the main glacial event within the ‘Devensian’ (Clark et al., 2004, 2012). During the Dimlington Stadial the Last British-Irish Ice Sheet extended across much of northern (Evans et al., 2009; Livingstone et al., 2012), northwest (Thomas et al., 2004; Livingstone et al., 2010; Chiverrell et al., 2016) and eastern England (Evans et al., 1995; Catt, 2007; Boston et al., 2010; Davies et al., 2012b; Roberts et al., 2013; Busfield et al., 2015; Dove et al., 2017). It resulted amongst other things in the development of extensive areas of ice-sculpted terrain (Fig. 9c), deposition of thick till and outwash sequences (Fig. 9d) and more locally, the development of meltwater-scoured valleys (Fig. 9e). Major ice-dispersal centres were located within the Lake District, Wales, Southern Uplands and the Scottish Highlands and these fed several ice streams that transferred ice rapidly to the marginal areas (Evans et al., 2005; Clark et al., 2012).

The British-Irish Ice Sheet (BIIS) reached its maximum extent at c.27 ka some eight thousand years before the ‘global Last Glacial Maximum’ (Clark et al., 2012). However, it is important to note that different sectors of the ice sheet reached their maxima at different times reflecting internal glaciological controls on ice dynamics. For example, the margins of the Irish Sea Ice Stream reached their maximum position at the Isles of Scilly approximately 26–25 ka before collapsing and retreating rapidly northwards (Scourse et al., 2006; Smedley et al., 2017). By contrast, the North Sea Ice Lobe appears to have reached its maximum extent much later around 17–15 ka BP (Bateman et al., 2015). Our understanding of the Last British-Irish Ice Sheet during the Dimlington Stadial is rapidly evolving under the current work of the BRITICE Chrono project and readers are directed to outputs from this project for further updates. During the Dimlington Stadial, large areas of central England remained un-glaciated. This caused drainage flowing from upland areas such as the Pennines and Peak District to become dammed by the coincident ice mass resulting in the formation of extensive icedammed lakes across large parts of eastern England (Bateman et al., 2008; Murton et al., 2009; Yorke et al., 2012; Evans et al., 2017). Beyond the ice margin, much of the landscape of central and southern England comprised arctic tundra with the widespread development of a thick zone of permafrost (Busby et al., 2015) and deposition of coversand and loess (Catt, 1978; Antoine et al., 2003; Murton and Lautridou, 2003; Bateman et al., 2014). 3.2.3.4. The Lateglacial (MIS 2-1 transition). The Lateglacial period follows the Dimlington Stadial and occurred against a backdrop of gradual climatic amelioration (Walker et al., 2003), isostatic adjustment and rising sea-levels (Shennan et al., 2006). Drainage systems within previously glaciated areas re-adjusted to nonglacial boundary conditions. Many of the larger river systems extended onto emergent shelf areas such as the English Channel and southern North Sea (Fig. 3f; Bridgland, 2010). In the southern North Sea, a vast low-relief periglacial landscape called Doggerland was revealed following deglaciation and became habitable for Palaeolithic hunter-gathers from about 14 ka (Housley et al., 1997; Coles, 1998; Gaffney et al., 2007; Wygal and Heidenreich, 2014). Major river systems from eastern England drained north and north-eastwards across Doggerland towards the North Sea coastline which at the time was probably located between eastern Scotland and southwest Norway (Coles, 1998, 2000; Gaffney et al., 2007). Climatically, the Lateglacial period in Britain is characterised by a distinctive climatostratigraphy that is replicated at other sites in northern Europe and within the Greenland ice core record (Mangerud et al., 1974; Björck et al., 1998; Mayle et al., 1999; Walker et al., 1994, 2003; Lowe et al., 2008). The interval is characterised by a period of rapid warming into the Windermere Interstadial (the British correlative of the continental Bølling/ Allerød), then an abrupt phase of cooling known as the Loch Lomond Stadial (the British correlative of the continental Younger Dryas) before rapid warming at the onset of the current interglacial, the Holocene (Fig. 7d; Mayle et al., 1999). This climatic history produces a distinct expression within the British landscape. Firstly, the largescale loss of ice from the British Isles during the Windermere Interstadial; the regrowth of a limited ice cap in Scotland and corrie/cirque/cwm glaciers in the English and Welsh uplands during the Loch Lomond Stadial (McDougall, 2001; Wilson, 2002; Golledge, 2010; Brown et al., 2012). These subsequently disappear during the warming that occurred at the onset of the current interglacial. Secondly, the occurrence, in sites outside of the margin of the Loch Lomond readvance ice masses, a tripartite sediment sequence, particularly in lake basins that were

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formed by Dimlington Stadial ice retreat/advance (Fig. 4; Lowe and Walker, 1977; Mayle et al., 1999; Palmer et al., 2015). The tripartite sequence (see Mayle et al., 1999; Candy et al., 2016b for a more recent discussion) records a lower unit, deposited during the Windermere Interstadial. It is characterised by evidence for climatic warming both in its sedimentology (organic and/or marl-rich) and palaeocology (re-establishment of birch/pine open woodland in lowland Britain or the presence of thermophilous coleopteran species) (Atkinson et al., 1987; Walker et al., 1993, 1994, 2003). This unit is then overlain by a mineralogenic unit, deposited during the Loch Lomond Stadial, indicative of increased landscape disturbance, which contains fossil assemblages dominated by arctic and alpine coleopteran species and the replacement of woodland species by Artemisia and Rumex species. The final unit in the tripartite sequence corresponds to the onset of the current interglacial and sees a return to organic/marl rich sediments and the onset of warm climate conditions with the progressive development of birch/pine and then deciduous woodland. The scale of these climatic oscillations was large. Coleopteran evidence from sites such as Llanilid (south Wales) and Gransmoor (north Yorkshire) indicate that the peak of interstadial warmth may have been as warm as the current interglacial (albeit extremely shortlived) with cooling characterised by a decline in mean annual temperatures of around 15  C over c.2 ka (Atkinson et al., 1987; Walker et al., 1993, 2003). Recent studies have shown that the Lateglacial in Britain and the North Atlantic/European region in general, is a period of greater complexity that originally envisaged. Stable isotope and chironomid studies demonstrate that the Windermere Interstadial was characterised by three distinct warming peaks, punctuated by two abrupt cooling events (Brooks and Birks, 2000; Brooks et al., 2012; Marshall et al., 2002; Whittington et al., 2015; Candy et al., 2016b). This is consistent with the climatic stratigraphy of Greenland Interstadial 1 (GI1) in the Greenland ice core record and also existing ideas of a cold event (Older Dryas) separating the Bølling/Allerød on the continent (Fig. 7d; Lowe et al., 2008). Furthermore, it is now apparent that there are at least two discrete phases to the Younger Dryas, an early cold/dry phase and a later warm/wet phase and this climatic stratigraphy has implications for the timing of the Loch Lomond glacial readvance (Bakke et al., 2009; Lane et al., 2013). Whilst abrupt events during MIS 3 seem to have had negligible expression in the fluvial systems of southern and central Britain, there is clear and well-documented evidence for fluvial response to the abrupt events of the Lateglacial. In river systems such as the Gipping (at Sproughton, Suffolk), Kennett (Berkshire) and the Upper Thames (Ashton Keynes), there is clear evidence for fluvial system response during this interval, through low-energy sedimentation during the interstadial and high-energy erosion and deposition during the stadial (Rose et al., 1980; Collins et al., 1996, 2006; Lewis et al., 2001). These studies show, however, that such events are affected by different lag times in the response of stream power and sediment supply so that an abrupt cooling event such as the Loch Lomond Stadial may have generated an initial phase of intense down cutting followed by rapid and large-scale aggradation (Rose et al., 1980). The abrupt climatic fluctuations of this interval also highlight the rapid response of early human populations to climate change with evidence, at sites such as Gough’s Cave in Cheddar Gorge for human recolonization very early in the interstadial almost synchronously with the onset of warming (Jacobi and Higham, 2009, 2011). The archaeological record of the interstadial is complex with distinct differences in the lithic technologies present in the early and late parts of the interstadial. This may reflect response to the climatic shifts outlined above or changes in the fauna present in the landscape as Britain became increasingly wooded during the course of the interstadial (Jacobi and Higham, 2011a, 2011b).

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3.2.4. Holocene The Holocene, the current interglacial, is characterised by ongoing landscape adjustment following deglaciation and driven by rapid warming, isostatic adjustment and rising sea-levels (Lambeck, 1995; Shennan et al., 2000, 2006; Shennan and Horton, 2002). Emergent shelf areas around England gradually became submerged, resulting in the rapid shortening of river catchments, lengthening of coastlines and ultimately, the separation of Britain from continental Europe. Rising sea-levels led to the progradation of coastlines south and westwards through the western part of the southern North Sea culminating with the submergence of Doggerland by approximately 7 ka (Coles, 2000; Gaffney et al., 2007). The significance of these drowned landscapes or ‘submerged forests’ in documenting landscape evolution of this time was initially summarised by Reid (1913) with a modern perspective provided by Bickett and Tizzard (2015). Overall, the literature available on the Holocene Epoch in England is extensive and, as with the case of the Late Devensian, the detail is beyond the scope of this paper. However, it is worth highlighting that the resolution with which the current interglacial can be studied allows a much better understanding of the complex interplay between factors that drive the landscape and geological record than is possible for earlier periods. In particular, the effect of progressive sea-level rise on landscape and surface processes at the onset of an interglacial are better understood for the Holocene than for any previous interglacial. The key factors that appear to drive landscape change are (1) climate (and abrupt climatic events) (Mayewski et al., 2004), (2) paraglacial processes (Ballantyne et al., 2014), (3) isostatic adjustment to deglaciation (Shennan et al., 2006), (4) eustatic sea-level rise (Lambeck and Cheppell, 2001) and (5) anthropogenic activity (see Zalasiewicz et al., this volume). In England, the climate of the current interglacial does not stand out as being exceptionally different from any pre-Holocene interglacial. In fact, with respect to the indicators that are routinely employed to characterise the warmth of previous interglacials it would appear relatively cool (see Candy et al., 2016a). By contrast, deposits of MIS 5e, 9, 11 and the ‘Cromerian Complex’ are rich in fossil flora and fauna that are indicative of climates that were warmer than the present day (Candy et al., 2010, 2016a). Where these fossil indicators occur, they are typically associated with the Holocene Thermal Maximum of northwest Europe (8–6 ka BP, the Atlantic of the Blytt-Sernander system). The early Holocene of Britain as well as the wider North Atlantic region, appears to be punctuated by a number of abrupt cooling events. The most clearly expressed of these is the 8.2 ka event, a short-lived (c.160 yr) cold interval, purported to be driven by an abrupt re-organisation of the North Atlantic circulation system in response to the final effects of glacial meltwater input (Holmes et al., 2011; Thornalley et al., 2011). Whilst this event is clearly seen in the oxygen isotopic signal of speleothems and lacustrine carbonates in Britain, it is unclear whether it had a significant impact on the British landscape (Marshall et al., 2007). At sites such as Hockham Mere (Norfolk), there is an increase in nonarboreal pollen at the expense of tree pollen, particularly in Corylus, at the approximate time of this event (Bennett, 1983). Further afield there is evidence in the Scottish islands for the loss of tree populations and burning in the landscape at c.8 ka (Edwards et al., 2007). It is important to highlight that Edwards et al. (2007) have argued that it is difficult to identify the impact of abrupt events because they are occurring in association with human modification of the landscape. This is, of course, the key characteristic of the Holocene that makes it distinct from the earlier Pleistocene interglacials, as the impact of humans becomes increasingly clear within the British geological record. Numerous researchers have highlighted the role that humans have played in the modification of natural sediment budgets within British river

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systems, leading to the accumulation of thick deposits of overbank sedimentation that have no obvious natural analogue in the British record of pre-Holocene interglacials (see Brown et al., 2012). The increasing role of human modification of the landscape can be clearly seen at Diss Mere (Norfolk), which has been cited as being the most appropriate stratotype for the current interglacial in Britain. This sequence contains a near continuous record of the current interglacial from its onset, with much of the record being annually laminated. This pattern of lamination breaks down from c. 2500 yrs BP onwards due to accelerated sedimentation patterns associated with human alteration of the landscape (Peglar et al., 1989). Whilst the impact of humans on the British landscape makes the current interglacial distinct from previous interglacials (and glacials), the other key difference in the geological record is the absence of the megafauna that characterised Britain during the entire Quaternary (Stuart, 1982). The landscape-forming influence of this fauna should not be overlooked. It is possible that there are periods in the British Pleistocene record, such as MIS 7, where the occurrence of open grassland under fully interglacial conditions was a function of the opening of the landscape by herds of large mammals (Murton et al., 2001; Candy and Schreve, 2007). The recolonization of Britain by humans during the current interglacial has been widely discussed elsewhere but the importance of the North Sea, offering pathways for migration across the area known as Doggerland prior to the attainment of maximum sea-level, is becoming increasingly clear (Pettitt and White, 2012; Bickett and Tizzard, 2015). The very rapid recolonization that occurs at sites such as Flixton and Star Carr (Vale of Pickering, north Yorkshire) has been used as an indication of human use of the central North Sea whilst it was still habitable and passable (Jacobi and Higham, 2011a, 2011b; Pettitt and White, 2012; Candy et al., 2015b; Palmer et al., 2015). 4. Dynamic evolution of the English landscape 4.1. Tectonic-landscape interactions The role of tectonics in driving Cenozoic landscape development across England is a controversial topic. Some have inferred the absence of Mesozoic cover rocks across parts of England to demonstrate that these areas remained largely emergent during the Mesozoic (e.g. Cope et al., 1992) and tectonically stable during the Cenozoic meaning that the general form of the modern landscape may be tens of millions years old (e.g. Murray, 1992; Hancock and Rawson, 2002). However, it is now widely accepted that the general form of the English landscape is largely a product of Cenozoic exhumation (Trotter, 1929; Green, 1986; Lewis et al., 1992), with major landscape elements having formed during the Mid-Miocene and Pliocene and refined during the Quaternary (Booth et al., 2015). Geological evidence for Late Cenozoic uplift and erosion includes the development of widespread planation surfaces (Robson, 1944; Westaway, 2009a, 2010), raised beaches (Bates et al., 1997; Westaway, 2009a), river terrace sequences (Maddy et al., 2000; Maddy, 2002; Bridgland, 2006, 2010) and the thermal history of crustal rocks (Bray et al., 2002; Hillis et al., 2008; Holford et al., 2009; Green et al., 2012). The precise tectonic drivers of Late Cenozoic landscape development remain complex, with in some areas, upto 2 km of exhumation having occurred (Westaway, 2009a, 2010; Green et al., 2012). The abundance of Alpine-related structures across the UK and adjacent shelf areas (Stoker, 2002) suggest that northwardsdirected crustal shortening has occurred throughout the entire intraplate region during the Cenozoic (Hillis et al., 2008). Alpinerelated tectonism appears to have been particularly active during the Palaeogene (but probably continuing into the Miocene) driving

basin iversion across southern England (e.g. Weald, Wessex and London basins), eastern England and the Southern North Sea (e.g. Cleveland Basin, Flamborough and the Sole Pit Inversion). Since the beginning of the Miocene, the regional maximum compressive stress (s1) direction across southern, central and eastern England and the Southern North Sea has been orientated roughly northwest to north-northwest. The minimum horizontal (s3) stress, by contrast, has been orientated northeast to east-northeast perpendicular to s1 (Kooi et al., 1991; Cloetingh et al., 2005; Rasmussen, 2009; Williams et al., 2015). To the north, the regional stress patterns are more complex. The relative effect of magmatic underplating caused by the Iceland Mantle Plume during the Palaeogene was also significant, affecting both southwest (Westaway, 2010) and southeast England (Gale and Lovell, this volume) and driving southeastward tilting of England (Cope, 1994; Thomson, 1995). The total amount of Late Cenozoic tectonically forced exhumation remains contentious (e.g. Cope and Bowen, 2015; Gale and Lovell, this volume). Nevertheless, exhumation caused the inversion of several Mesozoic basins across England and has been used to explain the surface distribution of bedrock geology across England and specifically the absence of Mesozoic cover (e.g. Chalk Group) across northern areas (Japsen, 1998; Huuse, 2002). However, where the eroded material was deposited remains unresolved (Cope and Bowen, 2015; Westaway, 2017). For the past 10 Ma (Late Miocene to present day), the lack of visible Alpine-related deformation in the North Sea Basin suggests that England has been subjected to a low stress regime without either strong extensional or compressive tectonic stresses. Indeed, in situ stress measurements from deep boreholes and wells indicate that current crustal stresses beneath much of England are intermediate (s2) (Evans and Brereton, 1990). This implies that crustal deformation is preferentially taken-up through strike-slip faulting and this assertion is supported by focal mechanism plots for recent UK earthquakes (Baptie, 2010). Therefore, neither Alpine compression nor mantle plume activity readily explain crustal uplift in England during the last 10 Ma. In the case of the latter, continental drift has led to Britain migrating away from the Iceland Mantle Plume (now situated beneath Iceland) and its influence on uplift had probably ceased by the end of the Palaeogene. Recent research has indicated that since the Pliocene, climate-driven denudational isostasy has variably been the principal long-term driver of landscape development (Watts et al., 2000, 2005; Westaway et al., 2002, 2006, 2015; Lane et al., 2008; Westaway, 2009a, 2009b, 2010; Farrant et al., 2014). Therefore, the Late Miocene-Pliocene interval records a period when direct tectonic influences on landscape evolution declined and climate-driven landscape change became dominant. Much of the present shaping of the English landscape owes its form to climate-driven processes that have controlled weathering rates, mechanisms of erosion, sediment transportation and deposition. Movement of surface materials from ‘source’ (areas of erosion) to ‘sink’ (areas of deposition) has in-turn driven denudational isostasy with some areas that have experienced crustal buoyancy (‘source’ areas where material has been eroded) and others subsidence (‘sink’ areas such as basins where material has been deposited). A classic example of denudational isostasy is the Crag Basin of East Anglia, where loadinduced subsidence effectively ceased following the Anglian Glaciation and the realignment of regional drainage and removal of sediment supply from the basin (Lee, 2015; Westaway, 2017). Several models advocate that vertical uplift and subsidence are primarily accommodated by ductile flow within the upper mantle and lower crust (e.g. Westaway et al., 2002; Westaway, 2009a, 2009b, 2017). However, a poorly-understood piece of the jigsaw is how the pre-existing basement structure beneath England may influence crustal response to an applied stress and in-turn influence earth surface processes. In general terms, the response

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of these structures to an applied stress depends on the rheological properties of the crust (which are a function of composition, thickness and temperature), the magnitude of stress, the principal stress vectors (s1–s3) and the geometry of the basement structure relative to these stress vectors (Twiss and Moores, 1992). Deep boreholes, seismic data, gravity and magnetic data reveal that beneath England the basement structure is complex, encompassing several accreted (Caledonian) tectonic elements that collectively comprise the Avalonia Terrane (Pharoah et al., 1995, 1996; Pharoah, 1999; Beamish et al., 2016). The rocks and structure of the Avalonia Terrane record the cyclic opening and closure of several ancient ocean basins (e.g. the Iapetus, Rheic and Tornquist oceans), with the continents of Gondwana and Laurentia eventually amalgamating during the Late Palaeozoic (Variscan Orogeny) forming the supercontinent of Pangaea (see Beamish et al., 2016 for an overview). Uplift and subsidence driven by denudational isostasy, at least in-part, may be accommodated by reactivation of pre-existing basement structures (sometimes with different senses of displacement) or through the development of new structures in overlying strata (Chadwick, 1986; Sibson, 1995; Cloetingh et al., 2005; Westaway et al., 2006; Westaway, 2010). Evidence for post-Palaeogene fault activity in England is however extremely limited. This is due in part to the present low-stress regime, but also difficulties in identifying evidence for faulting and an under-representation of faults on published geological maps (Aldiss, 2013). Nevertheless, a small but growing body of evidence indicates the influence of post-Palaeogene faulting and vertical ground motion on recent landscape development (Ellison et al., 2004; Bingley et al., 2008; de Freitas, 2009; Aldiss, 2013) and critically, highlights potential links to deep basement structure (Aldiss, 2013). 4.2. Climate-landscape interactions Climate has influenced landscape evolution throughout the Cenozoic. However, initial and step-wise amplification of the global climate signal (Milankovitch Cycles) through the PlioPleistocene has intensified its role in driving denudational isostasy and surface geological processes. Key tipping points, according to Westaway et al. (2002) and Westaway (2017), have occurred at c.3.1 Ma, c.2 Ma and c.0.9 Ma and correspond to a strengthening of the orbital-forcing climate signal, climatedriven landscape response and an increase in rates of denudational isostasy. For the c.2 Ma and c.0.9 Ma tipping points, this implies lags of c.0.7–0.6 Ma and c.0.3 Ma between the initial transition to ‘obliquity-forcing’ (c.2.7–2.6 Ma) and ‘eccentricityforcing’ (c.1.2 Ma, MPT) and a harmonised landscape response. The c.3.1 Ma ‘tipping point’ requires further consideration. The ‘tipping point’ was apparently “ . . . initiated following the deterioration in climate starting at c.3.1 Ma, which marked the start of upland glaciation elsewhere in Europe . . . ” (Westaway et al., 2002, p. 559). This implies that exhumation in other parts of Europe drove uplift across England, which seems somewhat implausible given the complex localised crustal response to an applied (or removed) vertical load (see Section 4.1). Additionally, extensive upland glaciation in Britain and adjoining Fennoscandia is only considered to have started at around c.2.7–2.5 Ma (Jansen et al., 2000; Sejrup et al., 2005; Böse et al., 2012; Lee et al., 2011, 2012; Thierens et al., 2012). The MPT ‘tipping point’ (c.0.9 Ma) appears to have been particularly significant in terms of denudational isostasy driving modern landscape development due to the transition to ‘eccentricity-forcing’ (100 ka) climate cycles and the amplification of earth surface processes. It enhanced weathering and erosion rates, leading to larger sediment budgets and an increased efficiency of

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systems (e.g. slopes, rivers, glaciers) in transferring sediment from ‘source’ to ‘sink’ (Bridgland, 1994, 2000; Bridgland et al., 2014; Rose, 2010). The impact of eccentricity cycles upon landscape evolution was especially pronounced during cold climate stages where reduced vegetation cover, strong seasonal discharge regimes and more active periglacial/hillslope processes drove highly-dynamic earth surface processes (Bridgland, 1994). The net effect was to amplify upland (‘source’ areas) uplift and lowland/ basinal (‘sink’ areas) subsidence rates giving rise to the apparent ‘doming’ over central England (Bridgland et al., 2014; Westaway, 2017). Over the past 0.5 Ma, the clarity of individual glacial/ interglacial cycles within the geological record has been marked. Not only has this enabled robust characterisation of specific MIS within the UK but also critically, direct comparison to other European sequences and global records (e.g. Candy and McClymont, 2013; Bridgland and Westaway, 2014). Evidence for more abrupt (sub-Milankovitch) climate changes can only be recognised during the Last Glacial-Interglacial Transition and early Holocene. This probably reflects the resolution of the geological record coupled with the increased accuracy and precision of geochronology during this time-interval to depict different climatic events. Within the landscape itself, weathering has played a crucial role in progressively degrading materials and altering their mechanical properties and behaviour (Forster et al., 1991) to a state where they can be efficiently recycled by landscape processes such as hillslopes, rivers and glaciers. In the case of most lowland river systems, these are long-term agents of landscape evolution, with cyclical phases of river incision and terrace aggradation driven by denudational isostasy (driving uplift and subsidence) and orbitalforcing (driving catchment dynamics and sediment availability) (Bridgland and Westaway, 2014). Other lowland river systems (e.g. the Great Ouse, Welland, Solent and Exe) appear to be more complicated with either local controls on long-term river evolution, or catchments that respond to abrupt climate changes occurring at sub-Milankovitch time-scales (Briant et al., 2004, 2006, 2012; Westaway et al., 2006; Gao et al., 2007; Brown et al., 2010). Glaciers are also prominent agents of landscape evolution and have been active periodically throughout the Pleistocene (Clark et al., 2004; Lee et al., 2011) and possibly parts of the Holocene (Harrison et al., 2015). Their relative influence on landscape evolution depends on the scale and frequency of glaciation, the rheology of the substrate (i.e. composition, water content, temperature) and amount of ice-bed traction (Lee and Phillips, 2013; Benn and Evans, 2014). During the Quaternary, glaciation has resulted in the over-deepening and reshaping of mountain valleys (e.g. Benn and Evans, 2014), the resetting and realignment of drainage (e.g. Gibbard, 1988; Bridgland, 1994; Rose, 2009), deformation of the substrate (e.g. Williams et al., 2001; Burke et al., 2009; Phillips et al., 2013) and the widespread erosion and redistribution of surface materials (e.g. Boulton, 1982; Clayton, 2000). 5. Conclusions  This paper reviews the evolution of the English landscape from the Miocene to the present day highlighting the relevant drivers of landscape evolution. Much of England has undergone exhumation (upto 2 km) during the Cenozoic resulting in widespread removal or erosion of Mesozoic cover rocks.  The primary driver of long-term landscape development throughout the Cenozoic is tectonics. During the Early to MidMiocene, major tectonic controls include north-directed Alpine compression and the regional response of the crust to earlier (Palaeogene) magmatic underplating. Rapid uplift and exhumation led to the development of generally terrestrial conditions and an extensive hiatus across much of England.

Please cite this article in press as: J.R. Lee, et al., The Neogene and Quaternary of England: landscape evolution, tectonics, climate change and their expression in the geological record, Proc. Geol. Assoc. (2017), https://doi.org/10.1016/j.pgeola.2017.10.003

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 Many of the upland landscapes of England originate from either the Mid-Late Miocene or Pliocene with significant reshaping during the Quaternary.  Over the past 10 Ma, England has not been subjected to either a strong compressional or extensional stress regime with most modern crustal movements (earthquakes) driving strike-slip rather than dip-slip crustal displacement. The interval between the Mid-Miocene and Pliocene therefore corresponds to a period when direct tectonic controls on landscape development were progressively declining.  Step-wise amplification of the global climate signal during the Plio-Pleistocene led to climate-driven denudational isostasy becoming the principal agent of landscape evolution influencing the types and dynamics of earth surface processes. The cause of this evolution of the global climate signal was the amplification of the orbital forcing (Milankovitch cycles) signal which appears to have triggered major changes in landscape response to climate change at specific threshold boundaries. These include the switch from precession to obliquity-driven orbital cycles (c.2.0 Ma) and obliquity to eccentricity-driven cycles (c.0.9 Ma) albeit with lags of c.0.6 Ma and c.0.3 Ma respectively. Since the Mid-Pleistocene Transition (MPT), global scale climate changes have driven high-magnitude changes in landscape especially during cold stages when the majority of geomorphological work occurs.  Changes in orbital forcing drove progressively enhanced weathering and erosion rates, and more efficient earth surface systems (e.g. hillslopes, rivers and glaciers) that redistribute sediment from upland areas (source) to lowland and basinal areas (sinks). It is suggested that vertical strain within the crust caused by denudational unloading (source) and loading (sinks) is partly accommodated by the reactivation of pre-existing geological structure within the Mesozoic cover and Palaeozoic basement.  Although there is increasing evidence in long-term climate archives for abrupt millennial/centennial scale climatic events, it is only during the LGIT and the early Holocene that clear evidence for these are found in the British geological record.  A combination of relative stratigraphic techniques coupled with absolute dating (where available), allows construction of a Neogene and Quaternary stratigraphic framework. In many cases, particularly over the past 0.5 Ma, this can be compared directly with other global and regional records (e.g. marine isotope stages and sub-stages) allowing the expression of specific glacial/interglacial cycles to be investigated.  This stratigraphic framework has allowed a chronological framework to be developed for the rich Palaeolithic record. Inturn, this has enabled the archaeology to be employed as a potential dating-tool where geochronology is absent. It has allowed the timing of the earliest phases of human occupation to be constrained and the complex patterns of occupation and abandonment to be more clearly understood.

Acknowledgements The authors wish to thank many individuals for the discussions including Don Aldiss, Brian Baptie, Mark Bateman, David Bridgland, Jonathan Ford, Dag Ottesen, Simon Price, Peter Riches, Jim Rose, Danielle Schreve, David Schofield, Chris Thomas and Mark Woods. Katie Whitbread is thanked for her comments on an earlier version of the manuscript. David Bridgland and Jim Rose are warmly acknowledged for their constructive comments as reviewers which have really helped to strengthen the paper. JRL and RH publish with the permission of the Executive Director of the British Geological Survey (NERC).

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